The Astrophysical Journal Letters, 785:L20 (4pp), 2014 April 20
C 2014.
doi:10.1088/2041-8205/785/2/L20
The American Astronomical Society. All rights reserved. Printed in the U.S.A.
ABIOTIC OXYGEN-DOMINATED ATMOSPHERES ON TERRESTRIAL HABITABLE ZONE PLANETS
Robin Wordsworth and Raymond Pierrehumbert
Department of the Geophysical Sciences, University of Chicago, 5734 South Ellis Avenue, Chicago, IL 60622, USA
Received 2014 January 23; accepted 2014 March 10; published 2014 April 1
ABSTRACT
Detection of life on other planets requires identification of biosignatures, i.e., observable planetary properties that
robustly indicate the presence of a biosphere. One of the most widely accepted biosignatures for an Earth-like planet
is an atmosphere where oxygen is a major constituent. Here we show that lifeless habitable zone terrestrial planets
around any star type may develop oxygen-dominated atmospheres as a result of water photolysis, because the
cold trap mechanism that protects H2 O on Earth is ineffective when the atmospheric inventory of non-condensing
gases (e.g., N2 , Ar) is low. Hence the spectral features of O2 and O3 alone cannot be regarded as robust signs of
extraterrestrial life.
Key words: astrobiology – planets and satellites: atmospheres – planets and satellites: physical evolution – planets
and satellites: terrestrial planets – planet–star interactions
Online-only material: color figure
2. DEPENDENCE OF THE COLD TRAP ON THE
NON-CONDENSABLE GAS INVENTORY
1. INTRODUCTION
The rapid growth of exoplanet discovery and characterization
over the last two decades has fueled hopes that in the relatively
near future, we may be able to observe the atmospheres of Earthlike planets spectroscopically. Such targets will be intrinsically
interesting for comparative planetology, but also for the major
reason that they may host life. To search for life on exoplanets by
observing their atmospheres, we must first decide on spectral
features that can be used as biosignatures. Despite extensive
theoretical study of various possibilities, detections of molecular
oxygen (O2 ) and its photochemical byproduct, ozone (O3 ),
are still generally regarded as important potential indicators of
Earth-like life on another planet (Segura et al. 2005; Kaltenegger
et al. 2010; Kasting et al. 2013; Snellen et al. 2013).
Various authors have investigated the idea that abiotic oxygen
production could lead to “false positives” for life (Selsis et al.
2002; Segura et al. 2007; Léger et al. 2011; Hu et al. 2012; Tian
et al. 2014). For example, it has recently been argued that the
build-up of O2 to levels of ∼2–3 × 10−3 molar concentration in
CO2 -rich atmospheres could occur for planets around M-class
stars, because of the elevated XUV/NUV ratios in these cases
(Tian et al. 2014). Extensive atmospheric O2 buildup due to
H2 O photolysis followed by H escape may also occur on planets
that enter a runaway greenhouse state (Ingersoll 1969; Kasting
1988; Leconte et al. 2013). However, because by definition the
runaway greenhouse only occurs on planets inside the inner
edge of the habitable zone, it should not lead to identification of
false positives for life.
For planets inside the habitable zone, it is commonly believed
that H2 O photolysis will always be strongly limited by coldtrapping of water vapor in the lower atmosphere. The purpose
of this note is to point out that a mechanism for O2 buildup to levels where it is the dominant atmospheric gas exists
for terrestrial1 planets in the habitable zone around any star
type. The reason for this is that the extent of H2 O cold-trapping
depends strongly on the amount of non-condensable gas in the
atmosphere.
Previously, we have shown that the degree to which a
condensing gas such as H2 O is transported to a planet’s
upper atmosphere is determined primarily by the dimensionless
number M = pv L/pn cp Ts , where L is the specific latent
heat of the condensing gas, cp is the specific heat capacity at
constant pressure of the non-condensing gas (or gas mixture), Ts
is temperature, pv and pn are respectively the partial pressures
of the condensing and non-condensing gases in the atmosphere,
= mv /mn is the molar mass ratio between the two gases,
and all values are defined at the surface. M is essentially the
ratio of the latent heat of the condensing gas (here, H2 O) to
the sensible heat of the non-condensing gas (primarily N2 on
Earth; Wordsworth & Pierrehumbert 2013). Values of M > 1
(M < 1) correspond in general to situations where the upper
atmosphere is moist (dry).
Figure 1 shows the surface temperature dividing the moist
and dry upper atmosphere regimes as a function of pn for
a pure N2 − H2 O mixture. As can be seen, on a planet
with 1 bar of N2 , a surface temperature of >340 K is required for a moist upper atmosphere, in rough agreement
with detailed radiative–convective calculations (Wordsworth &
Pierrehumbert 2013). However, the required surface temperature is a strong function of pn . For 0.1 bar only ∼295 K is required, while for 0.01 bar the value drops to ∼255 K. In general
there is no reason to expect that Earth’s atmospheric nitrogen
inventory is typical for a rocky planet: in the inner solar system
alone, the range of atmospheric N2 as a function of planetary
mass spans 3.3 times (Venus) to 6.6 × 10−4 times (Mars) that of
Earth. Delivery and removal of volatiles on terrestrial planets is
dependent on an array of complex, chaotic processes, so wide
variations in inventories should be expected (Raymond et al.
2006; Lichtenegger et al. 2010; Lammer et al. 2009).
3. ABIOTIC OXYGEN ON PLANETS WITH
PURE H2 O ATMOSPHERES
The O2 buildup mechanism can easily be understood
intuitively by a thought experiment involving a hypothetical planet with a pure H2 O composition (Figure 2).
1
Here we define “terrestrial” in the standard (broad) way as describing any
planet of low enough mass that it does not possess a dense hydrogen envelope.
1
The Astrophysical Journal Letters, 785:L20 (4pp), 2014 April 20
Wordsworth & Pierrehumbert
450
moist regime
T [K]
s
400
350
300
dry regime
250
−2
10
(a)
−1
0
10
10
1
10
log [p /bar]
10
n
Figure 1. Surface temperature defining the transition between moist and dry
upper atmosphere regimes as a function of the surface partial pressure of
the non-condensable atmospheric component. Here, the non-condensing and
condensing gases are N2 and H2 O, respectively. Results using O2 , Ar, or CO2
as the non-condensing gas are similar.
Lacking atmospheric N2 , Ar, and CO2 , such a planet will initially have a pure H2 O atmosphere, with the surface pressure
determined by the Clausius–Clayperon relation (Andrews 2010;
Pierrehumbert 2011). If the planet has the same orbit and incident stellar flux as present-day Earth, it will most likely be in a
snowball state (Budyko 1969). However, because H2 O cannot
be cold-trapped when it is the only gas in the atmosphere, it
will be photolyzed by XUV and UV radiation from the host star
(primarily via H2 O + hν → OH∗ + H∗ ). The resultant atomic
hydrogen will escape to space at a rate dependent on factors
such as the XUV energy input and the temperature of the thermosphere, and hence the atmosphere will oxidize.2
In the one-dimensional limit with no surface mass fluxes,
atmospheric O2 will build up on such a planet until pn is high
enough to cold-trap H2 O and reduce loss rates to negligible
values. In three dimensions, the initial atmospheric evolution
may depend on the planet’s orbit and sub-surface heat flux/
transport rate, because on a tidally locked, ice-covered planet
with pure H2 O atmosphere, conditions on the dark side could
be so cold that even O2 would condense. However, on a planet
with Earth-like rotation and obliquity, all regions of the planet
receive starlight at some point in the year, so once the surface O2
inventory passed a given threshold, buildup of an O2 atmosphere
would likely be inevitable. In addition, for any planet, transient
heating events such as meteorite impacts would be able to force
transitions to a stable state of high atmospheric pressure.3
What about more general scenarios? First, we can relax the
assumption of zero downward flux at the surface and consider
cases where the created O2 can be used to oxidize the interior.
Then, redox balance dictates that atmospheric oxygen levels
must build up until the loss of hydrogen to space is balanced by
the surface removal rate of oxidizing material. For example, if
an O2 removal rate4 of 5×109 molecules cm−2 s−1 at the surface
is balanced by diffusion-limited H2 O loss, given an escape
(b)
Figure 2. Schematic of possible evolutionary pathways for an initially waterdominated planet exposed to stellar XUV and UV. (a) H2 O photolysis causes
O2 and other oxidized products to build up on the planet’s surface regions of
low net instellation. (b) Once sufficient O2 has built up, the planet can transition
to a state where a stable O2 atmosphere is present and hydrogen escape to space
is balanced by oxidation of the interior.
(A color version of this figure is available in the online journal.)
rate Φ = bH2 O−O2 fH2 O (HO−1
− HH−1
), the molar concentration
2
2O
5
of H2 O at the cold trap must be 3×10−3 mol mol−1 under Earth
gravity. Here bH2 O−O2 is the binary diffusion coefficient of H2 O
in O2 (Marrero & Mason 1972), HO2 and HH2 O are respectively
the atmospheric scale heights of O2 and H2 O, and fH2 O is the
cold trap H2 O molar concentration.
The surface O2 partial pressure required to match this coldtrap concentration, which can be calculated by integrating the
moist adiabat equation (Ingersoll 1969) as in Wordsworth &
Pierrehumbert (2013), depends on both the surface and cold-trap
temperatures. In Earth’s present-day oxygen-rich atmosphere,
the cold trap occurs at a relatively high Tt ∼ 210 K, due
primarily to the warming effect of ultraviolet solar absorption by
O3 (Andrews 2010). Given Ts = 288 K and Tt = 210 K, fH2 O =
3 × 10−3 mol mol−1 requires a surface O2 partial pressure of
0.15 bar. For a snowball planet with Ts = 240 K, this would drop
to 0.022 bar. By comparison, for Ts = 288 K and Tt = 140 K,
0.025 bars is required.6 Because O2 build-up should lead to O3
formation and hence stratospheric heating, O2 partial pressures
2
We assume here, as in previous work, that the efficiency of H2 O photolysis
is not a limiting factor on the rate of hydrogen escape.
−1
3 The latent heat of sublimation of O (L
2
O2 = 213 kJ kg ) is only around
one-tenth that of H2 O (Lide 2000). Hence with only 25% energy conversion
efficiency, the kinetic energy of an impactor traveling at 10 km s−1 with
density 3 g cm−3 would be sufficient to sublimate a 1 bar atmosphere of O2 on
an Earth-size planet if its radius was 19.2 km.
4 The actual rate of interior oxidation of an H O world with an oxygen-rich
2
atmosphere is difficult to calculate. For comparison, the average rate of
3+
oxidation due to Fe subduction to the mantle on Earth over the last 4 Gyr was
estimated as (1.9–7.1) × 109 molecules O2 cm−2 s−1 in Catling et al. (2001).
The relationship between Φ and fH2 O depends weakly on the homopause
temperature via the scale heights and bH2 O−O2 . For simplicity, Th = 300 K is
used here.
6 The T = 140 K calculation may underestimate the required surface O
t
2
partial pressure, because effective blocking of H2 O photolysis also requires the
cold-trap altitude to be lower than that at which the atmospheric opacity in the
UV becomes less than unity.
5
2
The Astrophysical Journal Letters, 785:L20 (4pp), 2014 April 20
Wordsworth & Pierrehumbert
of at least a fraction of a bar appear plausible once the planet’s
atmosphere reaches a steady state.
−4
10
−3
10
4. ABIOTIC OXYGEN ON EARTH-LIKE PLANETS
p [bar]
How would things change on a more complex planet where
other atmospheric constituents were present? First, if the atmosphere contains some N2 or Ar, the amount of O2 required
to block H2 O escape will clearly be decreased, and increased
horizontal heat transport would reduce the likelihood of atmospheric bistability via O2 condensation in the planet’s regions of
low surface instellation. Reduced gases such as methane, which
can be outgassed from a planet’s interior by abiotic processes
(Levi et al. 2013; Guzmán-Marmolejo et al. 2013), could have
lifetimes similar to those on Earth today in an O2 -rich atmosphere, although variations in O3 and NOx concentrations as
a function of UV levels and atmospheric composition might
alter this (Wayne 2000). In addition, volcanically emitted sulphur species and heterogenous chemistry will also affect the
atmospheric redox balance. Future investigations using photochemistry models will allow constraints on the importance of
these effects as a function of the water loss rate.
Surface/interior redox exchanges are another source of complexity on a low-N2 Earth-like planet. If the planet forms with a
hydrogen envelope that is lost to space early on (e.g., Genda &
Ikoma 2008), its crust and oceans should initially be reducing, and the oxidized products of H2 O photolysis might react
rapidly with the surface at first. However, as long as this occurred, the upper atmosphere would remain H2 O-rich and rapid
photolysis could continue. Over time, the planetary surface and
interior would become oxidized, decreasing their ability to act
as an oxygen sink. Assuming Earth’s present-day XUV flux, a
lower limit on H2 escape from a hydrogen-rich homopause is
∼4 × 1010 molecules cm−2 s−1 (Tian et al. 2005). Given this,
an N2 -poor Earth could lose 2.1 × 1022 moles of H2 O over
4 Gy, or 28% of the current ocean volume.7 This translates to
66.2 bar of atmospheric O2 —a large enough quantity to cause
significant irreversible oxidation of the solid planet and hence a
strong decrease in the reducing power of the surface. Because
XUV fluxes are greatly enhanced around young dwarf stars in
general, total water loss could be many times this value in some
cases (Ribas et al. 2005; Ribas et al. 2010; Linsky et al. 2014).
Finally, an Earth-like planet could have CO2 outgassing,
plate tectonics and hence the potential for a carbonate–silicate
weathering feedback (Walker et al. 1981). The CO2 cycle on
an initially anoxic planet without N2 or Ar would be complex,
because CO2 condenses at relatively high temperatures (Lide
2000) but has low compressive strength in solid form (Clark &
Mullin 1976). In the absence of ocean/interior heat transport
processes, outgassed CO2 could build up on the low instellation
regions of a planet until the return flow of CO2 glaciers became
sufficient to transport it back to high instellation regions.
Setting aside the complexity of the full climate problem for
future study, we can nonetheless demonstrate the potential for
O2 build-up in cases where CO2 levels are such that the planet
has an Earth-like global mean surface temperature. Figure 3
shows the variation of atmospheric temperature and H2 O molar
concentration with atmospheric N2 content calculated using the
same methodology as in Wordsworth & Pierrehumbert (2013),
for an Earth-like planet at 1 AU around a Sun-like star, assuming
−2
10
−1
10
0
10
100
150
200
T [K]
250
300
−4
10
−3
p [bar]
10
−2
10
−1
10
0
10
−10
10
−5
10
fH O [mol/mol]
0
10
2
Figure 3. Atmospheric (a) temperature and (b) H2 O molar concentration
in thermal equilibrium as a function of pressure, as simulated by the onedimensional radiative–convective model. In each case the atmospheric composition is N2 − CO2 − H2 O. For the dotted, dashed and solid lines, the N2
inventories are 1, 0.17 and 0.007 times that of present-day Earth, and the dry
CO2 molar concentration is 1 × 10−3 , 0.1 and 0.9 mol mol−1 , respectively. As
can be seen, the upper atmosphere is moist when N2 levels are low, implying
rapid H2 O photolysis.
an N2 –CO2 –H2 O atmosphere with tropospheric H2 O relative
humidity of 0.5. In each case, the CO2 molar concentration has
been chosen to yield close to Ts = 288 K in equilibrium. As
can be seen, once the N2 content drops below a few percent of
that on present-day Earth, the high atmosphere becomes rich in
H2 O, implying rapid photolysis and hence planetary oxidation.
Hence we may conclude that even planets that are Earth-like
in all respects except for the N2 content of their atmospheres
have the potential to build up O2 abiotically until it is a major
atmospheric constituent.
5. CONCLUSION
Because O2 can become the dominant gas in the atmosphere
of a lifeless planet, alone it cannot be regarded as a robust
biosignature. Our results do not necessarily rule out its utility in
every case. However, they do demonstrate that the situation is
considerably more complex than has previously been believed,
with the likelihood of an abiotic O2 -rich atmosphere emerging
a complicated function of a planet’s accretion history, internal
chemistry, atmospheric dynamics and orbital state. Investigation
of the range of possibilities for terrestrial planets with variable
N2 and Ar inventories should be a rich area for future theoretical
research that will help to expand our understanding of climate
evolution mechanisms. Nonetheless, for a specific exoplanet,
7
In this calculation, we assume that 50% of the escaping hydrogen is
outgassed directly from the mantle.
3
The Astrophysical Journal Letters, 785:L20 (4pp), 2014 April 20
Wordsworth & Pierrehumbert
even detailed modeling might not lead to a definite conclusion
given the inherent uncertainties in processes such as volatile
delivery during formation.
Observationally, there may still be a way to distinguish the
scenarios we discuss here, but only if a reliable way is developed
to retrieve the ratio of O2 to N2 or Ar in an exoplanet’s
atmosphere. In principle this may be achieved by analyzing
the planet’s spectrally resolved phase curve (Selsis et al. 2011),
in transit by measurement of the spectral Rayleigh scattering
slope (Benneke & Seager 2012) in a clear-sky (i.e., aerosolfree) atmosphere, or possibly via spectroscopic observation of
oxygen dimer features (Misra et al. 2014). More work will be
required to assess the potential of these techniques to determine
O2 /N2 mixing ratios in realistic planetary atmospheres.
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R.W. acknowledges support from the National Science Foundation and NASA’s VPL program. This article benefited from
discussions with F. Tian, R. de Kok, S. Rugheimer and
D. Sasselov.
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